OXYGEN AND HYDROGEN ISOTOPES IN THE HYDROLOGIC CYCLE 1996.pdf · March 16, 1996 8:49 Annual Reviews...

March 16, 1996 8:49 Annual Reviews GAT-8 AR10-08

Annu. Rev. Earth Planet. Sci. 1996. 24:225–62Copyright c© 1996 by Annual Reviews Inc. All rights reserved

OXYGEN AND HYDROGEN ISOTOPESIN THE HYDROLOGIC CYCLE

J. R. Gat

Department of Environmental Sciences and Energy Research, WeizmannInstitute of Science, 76100 Rehovot, Israel

KEY WORDS: isotope fractionation, evapo-transpiration, precipitation, soilwaters,groundwaters

ABSTRACT

Changes of the isotopic composition of water within the water cycle providea recognizable signature, relating such water to the different phases of the cycle.The isotope fractionations that accompany the evaporation from the ocean andother surface waters and the reverse process of rain formation account for the mostnotable changes. As a result, meteoric waters are depleted in the heavy isotopicspecies of H and O relative to ocean waters, whereas waters in evaporative systemssuch as lakes, plants, and soilwaters are relatively enriched. During the passagethrough the aquifers, the isotope composition of water is essentially a conservativeproperty at ambient temperatures, but at elevated temperatures, interaction withthe rock matrix may perturb the isotope composition. These changes of the isotopecomposition in atmospheric waters, surface water, soil, and groundwaters, as wellas in the biosphere, are applied in the characterization of hydrological system aswell as indicators of paleo-climatological conditions in proxy materials in climaticarchives, such as ice, lake sediments, or organic materials.

1. INTRODUCTION

The basic tenet of the hydrologic cycle as expounded by the preacher, that“into the sea all the rivers go and yet the sea is never filled, and still to theirgoal the rivers go” (The Jerusalem Bible, English translation, 1970), remainsvalid until today. However, the simple concept of a steady-state one-cycle

2250084-6597/96/0515-0225$08.00


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Figure 1 The hydrologic cycle, in relative flux units (100 units equals the marine evaporationflux). Figures have been rounded up. (Adapted from Chow 1964.)

system needs to be qualified. As shown schematically in Figure 1, the marinepart of the cycle, namely the evaporation from the ocean and the return flux byatmospheric precipitation into the ocean, accounts for close to 90% of the annualflux. However, in the remaining continental part of the cycle, the precipitatedwater is recycled by evapo-transpiration a number of times before returning tothe oceans as river and groundwater runoff. Moreover, intermediate storagereservoirs such as glaciers and deep groundwater systems are changeable on avariety of time scales, in response to a changing climate.

Compared to the 1.35× 109 km3 of oceanic waters, the groundwater and icereservoirs on land are estimated at 82× 106 and 27.5× 106 km3, respectively.Atmospheric moisture constitutes only about 10−5 of the volume of the oceanwaters.

Measurements of the isotope composition of water in the various componentsof the water cycle has enabled the identification of different water masses andthe tracing of their interrelationships. Furthermore, measured variations (intime) of the isotopic composition (of proxy materials) in climatic archives such


O AND H IN THE HYDROLOGIC CYCLE 227

as in ice cores, lake and ocean sediments, and of plant material, have provenuseful for paleoclimate reconstructions.

The abundance of the isotopes2H (deuterium, also marked as D) and18O inocean waters is close to D/H= 155.95× 10−6 (Dewit et al 1980) and18O/O= 2005.2×10−6 (Baertschi 1976), respectively, where H and O without a massassignment stands for the natural assembly of all isotopic species. Anothernatural oxygen isotope,17O, occurs at an abundance of about 1/10 that of18O,but because its variations in the water cycle parallels that of the heavier isotope(at 1/2 the extent) it has not been measured as a rule.

Abundance variations of± 30% for deuterium and± 5% for 18O have beenmeasured in natural waters (Boato 1961). These variations are only to a neg-ligible degree the result of nuclear interactions that produce or consume theseisotopes; they rather result primarily from the fractionation of the isotopic wa-ter species in isotopic exchange reactions (equilibrium isotope effects) or inchemical or biochemical reactions involving the isotopic species (kinetic andvital effects), or as a result of different diffusion rates of the isotopic watermolecules (transport effects). In the water cycle, the most significant processin this respect is that of phase changes, from vapor to liquid or ice and viceversa.

The ocean, being the largest water reservoir and relatively homogeneous,was chosen as the reference standard for theδ scale of both oxygen and hydro-gen isotopes in water samples (Craig 1961a). Theδ-(permil) value is definedasδ‰ = (Rx/Rstd− 1) × 103, whereR is the atom ratio D/H and18O/16O,respectively. Positiveδ values thus signify an enrichment of the heavy iso-topic species relative to the standard (SMOW); negative values indicate theirdepletion. Because the variability inδ(D) is typically larger by almost an orderof magnitude than that ofδ(18O), the usual way of presenting isotope data ofwater samples is on aδ(D) vs δ(18O) diagram with theδ(D) scale compressedby a factor of ten.

Ever since the earliest measurement of isotopic abundance in the 1930s inEurope and Japan (Rankama 1954) and by the Chicago school of HC Urey, us-ing the Nier-McKinney double-inlet, double-collector mass-spectrometer (Nier1947, Mckinney et al 1950), it has been observed that freshwater samples aredepleted in the heavy isotopes, whereas the residue of evaporating surface wa-ters such as African lakes are enriched (Craig 1961b). A correlation between thedepletion of the deuterium and oxygen-18 in freshwaters was established; thelocus line of such samples inδ-space was defined asδ(2H) = 8δ(18O)+10(‰).This line is now known as the Global Meteoric Water Line (GMWL).


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2. ISOTOPE FRACTIONATION IN THE WATER CYCLE

2.1 Fractionation FactorsThe strength of chemical bonds involving different isotopes of an element differslightly, with the bond involving the lighter isotope usually weaker and easierto disintegrate. The rate of chemical reactions in which such a bond is brokenwill then show a (kinetic) isotope effect (Melander 1960). Similarly, and basedon the general thermodynamic principle that equilibrium systems tend towarda state of minimum energy, the heavy isotopes will favor that part of the sys-tem in which they are more strongly bound. This principle applies as well toequilibrium between different phases. These are quantum effects, which areappreciable at low temperatures and disappear at higher temperatures.

One defines a unit isotopic fractionation factor (α) in a system of two molec-ular species, A and B, with a common element or in two coexisting phases ofsome material as follows:

αA-B = RA

RB= (Ni /Nj )A

(Ni /Nj )B,

whereNi andNj are the amounts of the isotopes of the common element andR= Ni /Nj .

When the system is in isotopic equilibrium, one denotes the fractionationfactors as eitherα+ or α+, where by the convention introduced by Craig &Gordon (1965),α∗< 1; α+> 1 so thatα+ = 1/α+. The values of such anequilibrium fractionation factor depend only on the molecules and isotopesinvolved and are strongly temperature dependent, following a rule such as

lnα∗ = C1

T2+ C2

T+ C3.

Equilibrium fractionation factors can be calculated based on spectroscopic datausing the partition functions of the molecules concerned. A number of reviewson this subject have been published (Bigeleisen & Mayer 1947, Urey 1947,Boato 1961, Bigeleisen et al 1973).

In a more general manner, an isotopic fractionation factor can be defined fora system from which there is an outgoing flux with different isotope ratios thanin the bulk of the material:

αK = d Ni /d Nj

Ni /Nj.

One usesαK to distinguish this case from one where an equilibrium fractionationfactor applies.

Such a formulation applies when part of the system is removed by a chemicalreaction (in which caseαK can be called akineticfractionation factor), but it will



also apply to the case where material is removed by diffusion or outflow, e.g. byeffusion through an aperture. In the latter cases the termtransportfractionationfactor may be preferred. Unlike true kinetic fractionation factors, which likethe equilibrium factors are strongly temperature dependent, the transport factorsshow only slight (positive) temperature coefficients.

The values of the equilibrium fractionation factors (α+) for the water-vaporphase transition are given by Majoube (1971) as 1.0098 and 1.084 at 20◦C and1.0117 and 1.111 at 0◦C, respectively, for18O and D. For the ice-liquid transitionthe values ofα+ (at 0◦C) are 1.0035 for18O and 1.0208 for D (Arnason 1969).

2.2 The Rayleigh EquationsIf material is removed from a (mixed) system containingNi andNj molecules,respectively, of two isotopic species and if the fractionation accompanying theremoval process at any instance is described by the unit fractionation factorα, then the evolution of the isotopic composition in the remnant material isdescribed by the following equations, assumingN � Ni , whereNi +Nj = N,so that it follows thatN ' Nj (which holds for the natural isotope abundanceof the lighter elements, e.g. H, N, C, and O):

d R

d N= 1

N

(d Ni

d N− Ni

N

)= R

N(α − 1) (1a)

and

d(ln R)

d(ln N)= (α − 1). (1b)

Equation 1a can be immediately integrated from an initial condition(R0, N0) toany given stage, providedα does not change during the course of the process:

R= R0

(N

N0

)(α−1)

= R0 · f (α−1), (3)

where f = N/N0 is the fraction of the material remaining.When the isotopic species removed at every instant are inthermodynamics

equilibriumwith those remaining in the system, we have the circumstances ofthe so-called Rayleigh distillation. Under these conditions, the unit separationfactor will be a thermodynamic equilibrium constantα∗, i.e. the vapor pressureratio of the isotopic water molecules for the liquid to vapor transitions or theequilibrium constant of an isotopic exchange reaction.

The isotopic compositionR of the remaining material, as a function off ,when subject to Rayleigh fractionation, is shown in Figure 2. Obviously, theisotope composition of the material that is removed at every instance describes


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Figure 2 Isotopic change under Rayleigh, closed, and steady-state conditions for fractionationprocesses with a unit fractionation factor ofα = 1.01 for initial liquid composition ofδ = 0. A, B,C: Rayleigh conditions for residual water, continuously removed vapor, and total vapor removed,respectively; D, E: liquid and vapor for two coexisting phase systems; F: liquid approaching steadystate (time axis not in proportion).



a curve parallel to that of the remaining fraction. The integrated curve, givingthe isotopic composition of the accumulated material thus removed, is alsoshown. Material balance considerations require that the isotope content of thetotal accumulated amount of the removed material approachesR0 as f → 0.

2.3 Isotope Fractionation in Open, Closed,and Steady-State Systems

The system for which the Rayleigh equations apply is an open system fromwhich material is removed continuously under condition of a fixed fractiona-tion factor. However, other systems can be envisaged, to which the same unitfractionation factor applies but under other boundary conditions for the inte-gration of the unit process. One such system is the so-called closed system(or two-phase equilibrium model), first described by S Epstein (unpublishedlecture notes), where the material removed from one reservoir accumulates ina second reservoir in such a manner that isotopic equilibrium is maintainedthroughout the process. An example would be the condensation of vapor todroplets in a cloud.

If the timet = 0 all material is in reservoir #1, then

N2 = N1.0− N1, Ni2 = Ni1.0 − Ni1,

and also

R2

R1= α.

Defining f = N1N1,0

, then

Rf = R0

α − f (α − 1)(3a)

or in δ(‰) nomenclature

δ f = δ0− ε(1− f )

α − f (ε/103). (3b)

Unlike in the open system,δ f→0 −→ (δ0−ε)/α, limiting the degree of isotopeenrichment.

A third model is that of a throughflow system in which the outgoing (frac-tionated) flux is replaced by an inflow with isotope compositionδin. At steadystate, the isotope composition of in- and outflowing fluxes are obviously equal.If the outflowing flux is related to the reservoir’s isotope composition by a frac-tionation factorα, then the steady-state isotopic composition in the reservoir


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will be δss= (δin − ε)/α, while the approach to steady-state composition canbe expressed as (Zimmermann 1979)

δss− δδss− δ0

= exp(−αt/τ ), (4)

whereτ is the characteristic turnover time of the system, given by the volumeto flux ratio (τ = V/F).

Note that for the same unit fractionation factor the resultant evolution of theisotopic composition is quite different under these different circumstances, asshown in Figure 2.

In all examples given so far, a constant unit fractionation factor is assumedto apply throughout the process. This is not, however, always the case. Forexample, as will be discussed, the rainout from an air mass is usually the resultof a continuous cooling of the parcel of air concerned; the cooling results in achange of the unit fractionation factor for the vapor-to-water (or vapor-to-ice)transition during the evolution of the precipitating system. Although the differ-ential form of the Rayleigh equation (Equation 1a) still applies, the integrationobviously has to be carried out for a prescribed change ofα throughout theprocess [as was first demonstrated by Dansgaard (1964)].

Another conspicuous example of a changing “effective” fractionation factoris that of the evaporation of water from a surface water body into the atmosphere.In this case, the change is the result of the changing boundary conditions (ofthe isotopic gradient, primarily) rather than a change of the unit fractionationfactors themselves.

2.4 Isotope Fractionation Accompanying Evaporationof Surface Waters

Water-air interaction balances two opposing water fluxes, one upward from thesurface and the other a downward one of atmospheric moisture. At saturation,i.e. when atmospheric humidity is 100% in terms of the saturation vapor pres-sure at the liquid surface, this interaction will bring the liquid water and airhumidity into isotopic equilibrium with one another. Such a situation appar-ently occurs at cloud base between the falling rain droplets and the ascendingair.

When the air is undersaturated, a net evaporation flux results, in which therate-determining step is the diffusion of water vapor across the air boundarylayer in response to the humidity gradient between the surface and the fullyturbulent ambient air. Three factors are involved in determining the overallisotope fractionation: 1. the equilibrium isotope fractionation of the liquid tovapor phase transition, 2. a fractionation resulting from the diffusion across theair boundary layer, and 3. the back flux of the atmospheric moisture.



The most useful model for the isotope fractionation during evaporation,which has stood the test of time, is that of Craig & Gordon (1965). Themodel is based on the Langmuir linear-resistance model for evaporation andis schematically shown in Figure 3. The following assumptions underlie thismodel:

(a) equilibrium conditions at the air/water interface, so that the relative humidityis h = 1 andRV = α∗RL;

(b) a constant vertical flux, i.e. no divergence or convergence in the air column;and

(c) no isotopic fractionation during a fully turbulent transport.

The appropriate flux equations for the water substance (E) and isotopic molecul-es (Ei , either for HDO or H2

18O) are then

E = (1− h)/ρ; ρ = ρM + ρT

and

Ei = (α∗RL − h · Ra)/ρi ; ρi = ρi,M + ρi,T.

Figure 3 The Craig-Gordon evaporation model.


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Theρ terms are the appropriate resistances as shown in Figure 3; subscriptsM and T signify the diffusive and turbulent sublayer, respectively. Subscript arefers to a free atmosphere above the evaporating surface.

The isotope composition of the evaporation flux is

RE = Ei /E = (α∗RL − hRa)

(1− h) · ρi /ρ. (5a)

This can be translated intoδ units,δ = (R/Rstd− 1), to give

δE = α∗δL − hδa− ε∗ −1ε(1− h)+1ε/103

≈ δL − hδa− ε∗ −1ε(1− h)

, (5b)

where

ε∗ = (1− α∗) · 103

and

1ε = (1− h)

(ρi

ρ− 1

)× 103.

The term [(ρi /ρ) − 1] can be evaluated as follows: In the linear resistancemodelρi = ρi,M + ρi,T andρ = ρM + ρT;

ρi

ρ= ρM,i + ρi,T

ρM + ρT= ρM

ρ· ρM,i

ρM+ ρT

ρ· ρi,T

ρT,(

ρi

ρ− 1

)= ρM

ρ

(ρM,i

ρM− 1

)+ ρT

ρ

(ρi,T

ρT− 1

).

The second term on the right-hand side is eliminated on the assumption thatρi,T = ρT, so that

1ε

103= (1− h)

[ρM

ρ

(ρM,i

ρM− 1

)]. (6a)

For the case of a fully developed diffusion layerρM ∝ D−1m , where Dm is

the molecular diffusivities of water in air. For a rough interface under strong(turbulent) wind conditions, the transient eddy model of Brutsaert (1965) can beapplied, withρM ∝ D−1/2

m . At more moderate wind speeds a transition from theproportionality ofD−2/3 to D−1/2 can be expected (Merlivat & Contiac 1975).

The ratio of the molecular diffusivities in air, of the pairs H218O/H2

16O and1HDO/1H2O, has been determined by Merlivat (1978) as 0.9723 and 0.9755,respectively. Note that these values differ from those calculated by simple gaskinetic theory, based on the molecular weights only, apparently due to somehydrogen bonding in the gas phase. Accordingly, the1ε-term is expressed as

1ε = (1− h) · θ · n · CD · 103, (6b)



Figure 4 The marine evaporation process in terms of theδ-diagram.

where 0.5 ≤ n ≤ 1;CD(18O) = 28.5.5‰ andCD(D) = 25.115‰. Theweighting term,

θ = ρM

ρ= (1− h′)

1− h,

(see Figure 3) can be assumed equal to 1 for a small body of water whoseevaporation flux does not perturb the ambient moisture significantly (Gat 1995);however,θ has been shown to have a value of 0.88 for the North American GreatLakes (Gat et al 1994) and a value of close to 0.5 for evaporation in the easternMediterranean Sea (Gat et al 1995). For an open water body under naturalconditions, a value ofn = 1/2 seems appropriate. In contrast, in the case ofevaporation of water through a stagnant air layer, such as from the soils (Barnes& Allison 1988) or leaves (Allison et al 1985), a value ofn ∼ 1 fits the data well.

Figure 4 shows in schematic manner theδ(D)− δ(18O) relationships in thisprocess.δE andδL define a line inδ(18O) vsδ(D) space, called the evaporationline (EL) whose slope is given by

SE(δ) = [h(δa− δL)+ ε]2H

[n(δa− δL)+ ε]18O

. (7)


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As shown in Figure 4, from material balance considerations, both the initialwater composition, the evaporated moisture, and the remnant liquid (such aslake waters, surface ocean water, or soil waters) all lie on this line. Obviouslythe humidity, the isotope composition of the air humidity, and the fractionationfactorsε∗ and1ε [which in turn are temperature (forε∗) and mechanism (for1ε) dependent] determine the slope. A special case is presented whenδL andδa are in isotopic equilibrium with each other for both isotopic species (H andO), in which caseSE = (ε∗ + θnCD)2H/(ε

∗ + θnCD)18O, independent ofh.This slope isS= 3.54 whenθ = 1 andn = 1/2 and is lower for the case ofthe fully developed boundary layer, when forn = 1 the equilibrium slope canbe calculated to beSE = 2.69. Havingθ < 1 will increase the slope of therelevant evaporation line, as will be the case whereδL − δa > −ε∗.

The Craig-Gordon model does not account for the case where evaporation ofdroplets and spray contribute to the evaporation flux. Furthermore, if surfacewaters are not well mixed so that a concentration gradient can build up betweenthe evaporating surface skin and the bulk water, this needs to be taken intoaccount by introducing a resistance in the liquid into the linear evaporationmodel (Craig & Gordon 1965).

3. THE ISOTOPIC COMPOSITION OF ATMOSPHERICMOISTURE AND OF PRECIPITATION

To a first rough approximation, as shown in Figure 1, the water cycle can bedescribed by its marine part, which accounts for 90% of the water flux. Craig& Gordon (1965) in their first attempt to describe the isotopes of the ocean-atmosphere system indeed modeled the marine atmosphere as a closed cycle(Figure 5a) assumingδE = δp (whereδp, which is close toδ(18O) = −4‰andδ(D) = −22‰, is the worldwide average of precipitation) and that 2/3of the moisture in the ascending are rained out, resulting in air with a meanhumidity of 75%. The marine moisture’s isotopic composition is neither that ofthe evaporative flux(δE), nor is it in equilibrium with the ocean surface water;in the model it is assumed to be in isotopic equilibrium with the precipitation.

The isotope data of marine precipitation, collected by the IAEA-WMO pre-cipitation network from island stations and weatherships (an excerpt of whichis shown in Figure 6), show that marine air is not uniform and that a mean iso-topic value for marine precipitation isδ(18O) = −2.5 to−3.0‰, with variableδD values, rather than that assumed in the above-mentioned model. The morerealistic, though still grossly simplified model, shown in Figure 5b, assumesa value ofδp(

18O) = −3‰; δ(D) = −14‰ for marine precipitation, whichaccounts for 90% of the evaporative flux. Solving Equation 5b for both the



isotopes and assuming marine air in equilibrium with the marine precipitationat a temperature of 20◦C (with the humidityh and the factorθ as open param-eters) yield very reasonable values ofh = 0.743 andθ = 0.48. This value ofθ ∼ 0.5 should be noted, as one encounters this value repeatedly over largeevaporative systems (Gat et al 1995) as noted above.

As the marine air then moves over the coast and into the continents, the differ-ent marine air parcels appear to mix and homogenize, resulting in precipitationthat is closely aligned along the so-called meteoric water lines (MWL). Theseare the lines inδ-space whose formula isδ(D) = 8δ(18O)+ d, where d has beennamed the “deuterium excess” parameter by Dansgaard (1964); obviously theGMWL (Craig 1961b) is one such meteoric water line with d= 10‰. Thefurther removed from the vapor source, the more depleted the isotopic valuesof the precipitation on each of these lines. Dansgaard (1964) recognized fourparameters that determine this depletion in the isotopic values, including analtitude effect, a distance from the coast, and a latitude effect. All of theseare basically related to the wringing out of moisture from the atmosphere as aresult of cooling of the air mass, and indeed the correlation with the temperatureappears as the overriding factor (Yurtsever 1975). Only the fourth parameterdescribed by Dansgaard, the “amount effect,” has a more complicated structure(Gat 1980). Figure 7 shows the worldwide distribution of the oxygen-18 valuesin annual precipitation.

To explain these findings of terms of the Rayleigh equations described above,the early concepts on the evolution of the isotopic composition of “meteoric”waters envisaged an air mass (with water vapor mass N and a given isotopiccompositionRv) from which precipitation is derived by equilibrium condensa-tion, so that at any time:Rp = α∗Rv. The system then follows a “Rayleighlaw,” according to Equation 1b, with the fractionation factor ofα∗. In integratedform λ = λ0 + ε∗ ln f , with λ ≡ lnR; (α∗ − 1) ≡ ε∗ and f is the fraction ofremaining water vapor in the system( f = N/N0).

This equation establishes a linear relationship inλD-λ18 space under isother-mal conditions, i.e. whenα∗ remains fixed throughout the process. The slopesof such Rayleigh lines inλ-space vary from the value ofSλ = ε∗D/ε∗18 = 8.22to a value ofSλ = 9.6 over the temperature range of 30◦C to 0◦C, accordingthe fractionation factors as given by Majoube (1971).

Theλ- andδ(‰)-space are related by the equations:

R

Rstd=(

1+ δ

103

)and

dλ = d[ln(1+ δ/103)] = dδ/103

(1+ δ/103).


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(a)

Figure 5 (a) The marine-atmosphere isotope model of Craig & Gordon (1965). The first of theδ

values refers toδ(18O), while the value in parenthesis refers to the value ofδ(D).

Translation of the Rayleigh lines intoδ(‰)-space then imposes a deviationfrom linearity as the system becomes further removed from the zero of theδ-scale. Inδ units the Rayleigh law is expressed in the following form:

dδ/103

(1+ δ/103)= ε∗d ln f

and the Rayleigh slope inδ-space will then be given by

(Sδ)Rayl. = ε∗D(1+ δV/103)D

ε∗18(1+ δV/103)18O,

whereδ values refer to the vapor phase.This slope is not a constant, but changes with the magnitude ofδ(18O) and

δ(D). The correction term(1+ δV/103) becomes significant as the system isremoved from its original state and acquires large negativeδ values, as is thecase especially for the hydrogen isotopes. It is often neglected in other cases.



(b)

Figure 5 (b) Adaptation of the Craig-Gordon marine model to account for the terrestrial part ofthe hydrologic cycle.∞ indicates isotopic equilibrium between the vapor and precipitation fluxes.


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Figure 6 Monthly precipitation values at selected marine stations, based on data of the IAEAnetwork on aδ(D) vs δ(18O) diagram (Rosanski et al 1992).



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An isothermal situation, however, is not a realistic model for continuousrainout as required by the Rayleigh formulation. The only close-to-isothermalnatural system is the tropical rain forest, whereT ∼ 26◦C. In this case theabundant rain feeds on evapo-transpired (recycled) moisture so that the Rayleighplot does not adequately describe the isotopic relationship (Salati et al 1979,Gat & Matsui 1991).

A more generally realistic hydro-meteorological model is that of an air-mass that loses part of its water content during a continental passage as aresult of cooling by an orographic or latitudinal effect (Dansgaard 1953). Theisotopic evolution of the precipitation derived from such an air mass can thenbe calculated under the assumption of vapor/liquid equilibrium during rainout,which, however, occurs at progressively lower temperatures. In this case theisotope composition of the precipitation closely follows the linear relationshipof the MWL, as a consequence of the opposing effects of the(1 + δ/103)

correction term (which increases as the isotopic values are more depleted) andthe increased isotopic fractionation factor (theε∗ term), as a result of the lowertemperatures, as Taylor (1972) has shown.

The additive parameter in the meteoric water lines (which was defined as“d”) is evidently inherited from the air masses’ initial isotopic composition,as determined by the air-sea interaction regime: It remains relatively invariantduring the Rayleigh rainout. Different air-sea interaction conditions at thesource result in a series of parallel meteoric water lines (each of slope 8), forwhich the d-parameter in each case is related to the source of moisture (Gat1981), in particular to the moisture deficit above the interface at the site wherethe air-masses’ water content is acquired (Merlivat & Contiac 1975). Particularregional lines relating to some locally effective moisture sources have beenrecognized, such as the eastern Mediterranean–MWL of d∼ 20‰ (Gat &Carmi 1970) and a western Australian line with d∼ 14‰, etc.

The meteoric water lines discussed so far, which in essence are “Rayleighlines,” obviously describe a spatially distributed data set, which describes theevolution of a particular precipitating air mass. They also apply to different airmasses, with similar initial properties (water content, temperature, andδv,0),formed under comparable circumstances.

The physical reality behind the Rayleigh formulation of the rainout processis the isotopic exchange between the falling droplets and the ascending air in thecloud (Bolin 1959, Friedman et al 1962), resulting in precipitation that essen-tially “forgets” the isotopic label of very depleted isotopic values imprinted bythe in-cloud processes, establishing isotopic equilibrium with the ambient air.Indeed, to a good approximation the depletion in isotopic composition in precip-itation correlates well with the near-ground temperature, as shown by Dansgaard



(1964) and Yurtsever (1975), or more precisely with the temperature at the cloudbase (Rindsberger & Magaritz 1983). This simple scenario, which generallyfits the data rather well, is modified under exceptional circumstances, namely,

• in the case where snow or hail reaches the ground; the isotopic exchangethen does not occur, with the result that the precipitation is more depletedthan in the equilibrium situation. Often, in addition, the solid precipitationforms show higher d-values due to nonequilibrium condensation during thegrowth of ice particles (Jouzel & Merlivat 1984).

• in the case of precipitation from strongly convective systems [thunder clouds,cold fronts, and tropical clouds associated with the ITCZ (intertropical con-vergence zone)] which are characterized by strong local downdrafts. As aresult the raindrops do not interact with an averaged sample of the ambientair but only with a portion of the in-cloud air.

In the two cases cited above, the isotopic value in the precipitation is moredepleted than the true equilibrium precipitation. These situations are thenless effective in isotopic fractionation during rainout than the Rayleigh pro-cess proper, moderating the extent of isotopic depletion as a function of therainout.

An additional process that can affect the isotopic composition of precipitationand of the atmospheric moisture is the evaporation from falling rain dropletsbeneath the cloud base or from surface waters. Evaporation from the fallingrain results in the enrichment of the heavy isotopic species in the remnant dropalong the so-called evaporation line; since these usually have a slope of less than8 in δ(D) vs δ(18O) space, the resulting precipitation shows a smaller d-valuethan at the cloud base. On the other hand, as shown in Figure 8, the evaporatedmoisture’s isotope composition is characterized by larger d-excess value, so thatthe precipitation derived from an air mass into which the reevaporated moistureis admixed is also characterized by a large d-excess (Dansgaard 1964). Asimilar effect occurs with the evaporated flux from surface waters (Goodfriendet al 1989, Gat et al 1994).

The d-excess parameter has been shown to be a diagnostic tool for measuringthe contribution of evaporated moisture to the downwind atmosphere (Gat et al1994). The scaling factor for judging a change in the d-value of the perturbedatmosphere is given by

(dE− da) = (dL − da)

1− h+ (CD(D) − CD(18O)

)× θ × n,

with subscripts E, a, and L standing for the evaporation flux, the ambient mois-ture, and the liquid, respectively.


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Figure 8 Schematics of the addition of evaporated moisture from surface water into the ambientatmosphere on theδ-diagram. δp = δ value of precipitation;δw = δ value of the residual waterafter evaporation;δE = the evaporation flux;δa andδ′a are the atmospheric moisture before andafter mixing withδE.

Transpired moisture, in contrast, is usually not fractionated relative to soilmoisture, so that if the latter is derived from precipitation, the transpired fluxdoes not perturb the d-value of the atmospheric moisture (Salati et al 1979).

Some authors have attempted to fit a linear relationship (inδ-space) to pre-cipitation data from a particular station (usually the monthly averaged valuesas reported for the IAEA sampling network; IAEA 1992), naming such lines“local meteoric water lines.” Such a procedure is not always meaningful, forinstance when the various precipitation events are related to air masses with dif-ferent source characteristics or to different synoptic patterns, especially duringdifferent seasons.

3.1 Seasonal and Short-Term Variation in the IsotopicComposition of Precipitation

As seasons change, the isotopic composition of rain at any site will reflecttwo basic changes in the rain pattern. The first is the change of the sourcecharacteristics over the ocean due to the seasonal change in ocean temperature



and air-sea interaction conditions. This affects the d-value primarily. Further-more, different seasons have different degrees of rainout as dictated by thetemperature along the air mass trajectory (which fixes the position along the“Rayleigh” line). In addition, the deviation from the Rayleigh relationship byeither evaporative enrichment from falling raindrops, the admixture of reevapo-rated moisture, or the effects of snow and hail are also different for the differentseasons.

A typical seasonal march of precipitation values is shown in Figure 9, basedon data provided by the IAEA precipitation network (IAEA 1992). Both theheavy isotope content and the d-parameter vary over a yearly cycle, with thed-values for the winter precipitation usually somewhat higher. Evidently, the

Figure 9 Seasonal change ofδ(18O) and d for selected continental stations, based on IAEAnetwork data (IAEA 1992).


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effect of high d-values in snow (Jouzel & Merlivat 1984) and the reduceddegree of evaporation from falling rain droplets in winter when compared tothe summer conditions usually override the effect of higher humidity at theoceanic source areas, which would result in the opposite effect (Merlivat &Jouzel 1979).

There are still insufficient data to fully characterize shorter term isotopicchanges during a storm or in between individual precipitation events. It ap-pears that such variations span a wide range ofδ-values, up to± 10‰ in δ18Oat times. Some early studies by Epstein (1956), Bleeker et al (1966), and Mat-suo & Friedman (1967) indicated differences between rain produced by warmfront uplifting or cold fronts and associated thunder clouds. These studiesalso showed that the beginning part of most rainshowers are more enriched inthe heavy isotopes due to partial evaporation during the fall of rain dropletsthrough the air column. Later studies (Leguy et al 1983, Gedzelman et al 1989,Rindsberger et al 1990, Pionke & Dewalle 1992, McDonnel et al 1990) makeit appear that these differences reflect mainly the source of moisture and itsrainout history, and only to a lesser extent the local rain intensity (EM Adar,unpublished data from the Negev). However, a notable exception to this rule isgiven by very strong tropical rains, associated with the ITCZ and its toweringclouds, when precipitation with extremely depleted isotopic values are foundat the peak of the downpour (Matsui et al 1983).

It is also of interest that in a continental setting the average air moisture andthe daily precipitation are close to isotopic equilibrium with each another (Craig& Horibe 1967, Jacob & Sonntag 1991). This is not strictly true in a coastalsetting or in regions of climate transition (Matsui et al 1983).

4. PRECIPITATION-INFILTRATION-RUNOFFRELATIONSHIPS OVER LAND

4.1 Surface Interception of PrecipitationAs rain falls on the ground, a number of processes take place that modify theisotope composition of the incoming precipitation. Details of these processesdepend on the nature of the terrain and on the characteristics of the rain, suchas its amount and duration, intensity, and intermittency.

As shown in Figure 10 [which is an adaptation of an early model of Gat &Tzur (1967)], part of the incoming precipitation is intercepted on the canopyof the vegetation cover and in part lost by evaporation. Up to 35% of theprecipitation is thus lost in tropical forests (Molion 1987) and 14.2 and 20.3%,respectively, on deciduous and coniferous trees in the Appalachian Mountainsin the USA (Kendall 1993). While it evaporates, the water on the canopy,



Figure 10 The interception reservoirs during the precipitation to runoff transition, based on Gat& Tzur (1967).

as well as in other surface reservoirs, is enriched in the heavy isotopes. If afurther rainshower then follows before the canopy is dried up, this can flush theenriched residual of the partially evaporated waters to the ground.

In an attempt to recognize the isotopic signature imposed by the canopyinterception, the isotopic composition of the throughflow under a tree wascompared to that of the free-falling precipitation, as collected in a clearing of theforest (Leopoldo 1981, Saxena 1986, Dewalle & Swistock 1994). The resultsshow that the expected isotopic enrichment in the throughflow is overshadowedby selection of part of the rainwater during the process. Apparently the less-


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intense parts of the rainshower are preferentially lost. Since these usuallyconstitute the isotopically most enriched part of the rain [corresponding to theamount effect of Dansgaard (1964)], this selection process introduces a negativebias that cancels the signature of the evaporation process.

The isotopic change on an interception volumeV—whereaV is the firstrainshower,f is the fraction ofV remaining after evaporative water loss, andbV is the follow-up shower—that results in throughflow whenb > (1− f ) canbe modeled with some simplifying assumptions. The isotopic composition ofthe throughfall is then

δTh = f δA + (b− 1)δB

(b+ f − 1)+ f

∫ε′( f )d ln f

(b+ f − 1), (8)

where

dδ f

d ln f= h(δ f − δatm)− ε

(1− h)= ε′( f ), (9)

andδA andδB are the isotopic composition of the first and second rain events,respectively.

Whereas the first term is simply the weighted mean of the contribution ofthe first and second rainy spell, it is the second term that expresses the changein isotope composition due to evaporation. As was shown in a similar modelapplied to evaporation from the soil column (Gat 1987), this term has a max-imum for f ∼ 0.5 and vanishes asf approaches either 0 or 1. Although theeffect of evaporation from the canopy is thus difficult to quantify based on thelocal recharge flux, it is reflected in the downwind atmosphere in the form of anincrease of the d-excess parameter of the atmospheric moisture. Gat & Matsui(1991) thus interpreted an increase of close to 3‰ in the value of d over theAmazon Basin as the cumulative effect of reevaporation from the canopy ofthe rainforest. The transpiration flux, while larger than that of the interceptionflux, is not expected to produce a change in d.

In the semi-arid zone the plant cover is less developed. Here the directevaporation from temporary surface water impoundments that appear followingstronger rains are expected to produce the largest isotopic signal. Gat & Tzur(1967) calculated the change between the isotopic composition of infiltratingwater from such surface impoundments and the precipitation. For the climaticconditions of Israel the enrichment is less than 1‰ [inδ(18O)] on soils withinfiltration capacity of< 2 mm/hr in winter and∼ 2‰ in summer. On heavysoils, higher distortions of up to 5‰ are expected. The effect is minimal onsandy soils where one does not encounter free surface waters. The effect ofevaporation from within the soil column, and especially from well-ventilatedsand, is discussed below.



4.2 Surface RunoffAny water in excess of the holdup capacity of the surface reservoirs and of theinfiltrability of the soil appears as surface runoff. Stable isotopes have playeda decisive role in understanding the mechanism of their formation. A Cana-dian group pioneered this work (Fritz et al 1976). One chooses a particularrainfall event with extreme isotopic values and traces this pulse into the lo-cal and regional runoff. In the case of the temperature (wet) environment ofCanada Sklash et al (1976) found that the antecedent wetness in the topsoilwas involved in the runoff and that the last rainfall (which triggered the runoffevent) was diluted by the past rainfalls of the season. Some degree of evapo-rative enrichment, presumably from waters in the topsoil, was detected (Fritzet al 1976). Similar results were reported from other localities (Mook et al1974).

The situation is dramatically different in the arid zone. Here the deficiencyof rainfall and long intervals between rain events result in a dry top layer,the absence of a continuous plant cover, and bare rock surfaces denuded ofsoils. Loess areas are baked dry to form a rather impermeable “desert pave-ment.” Small amounts of rain that just wet the surface are then lost by completeevaporation. But even moderate rainfall events of just a few millimeters ofprecipitation result in local surface runoff from the hillsides. These combine,when the rain event is more continuous, to produce the phenomenal desert flashfloods on a regional scale. Isotopic studies in the highlands of the Negev Deserthave shown that the local runoff assumes the mean isotopic composition ofthe particular shower (Gat 1987) but that the large flash floods are consistentlydepleted in the heavy isotopes relative to the local rain (Levin et al 1980) whichgives rise to them. This effect has been attributed to the intra-storm isotopevariability where the most intense part of the shower shows depleted isotopicvalues (E Adar, unpublished data). Ehhalt et al (1963) have described a similareffect of depleted isotopic values in river runoff in South Africa.

The exception to these interrelationships are the sand dune areas in the dryzone, on which almost no surface flows occur because all incoming precipitationinfiltrates into the open pore structure of the sand. The subsequent evaporationfrom within the sand column is another story.

4.3 Lakes and Other Surface WatersThe enrichment of the heavy isotopic species in residual surface waters as a re-sult of the isotopic fractionation that accompanies evaporation was recognizedlong ago (Craig 1961b). This buildup of the isotopic enrichment and the recog-nizable isotopic signature of lake water is being utilized to establish the waterbalance of the lake, especially the ratio of inflow to evaporation fluxes, and as


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a tool for studying mixing in the lake. It is also a useful tracer for detecting theaddition of lake waters to adjacent groundwaters (Krabbenhoft et al 1994).

Under hydrologic steady-state conditions the isotopic buildup in a lake isgiven by

1δ = δL,SS− δin = (δa− δin + ε/h)[1+ Fin

E · (1−h)h

] , (10)

where Fin and E are the influx and evaporation terms, respectively. Evidently,changes in eitherδin, δa, the humidity (h), or the hydrologic balance of thelake are determining parameters. Higher enrichments up to a limiting value ofδL = δa+ ε/h are possible for situations such as a string of lakes, i.e. when anumber of evaporative elements act in series (Gat & Bowser 1991).

The subject of lakes, freshwater and saline ones, has been extensively re-viewed over the years. Gat (1995) summarizes the state of the art in thisrespect.

4.4 SoilwatersWaters infiltrating from the surface drain through the void spaces in the soil,once the “field capacity,” i.e. the amount of water held against gravity onthe grains of the soil by adsorption and capillary forces, is filled up. In veryhomogeneous soils one can recognize a piston flow pattern where the infiltrate ofeach season pushes the previous one ahead, as visualized by the tritium contentof the moisture column (Dincer et al 1974, Gvirtzman & Magaritz 1990). Inheterogeneous soils and fractured matrixes the pattern is more complicated.

The holdup in the soil column and transport through it does not, in itself, affectthe isotopic composition of the infiltrating waters, except as far as mixing be-tween moving and standing water parcels occurs. However, when evaporationfrom within the topsoil occurs during dry intervals between rain events, this re-sults in an enrichment of the heavy isotopes in the residual waters (Zimmermanet al 1967). At some depth beneath the surface an evaporating front develops:Above it water transport is predominantly in the gaseous phase through the air-filled intergranular space; below it laminar flow and molecular diffusion in thewater-filled pore space dominate. An isotopic concentration profile develops,representing a balance between the upward convective flux and the downwarddiffusion of the evaporative signature (Barnes & Allison 1988). Isotopic en-richments of18O in excess of 10‰ relative to the precipitation are found indesert soils. Because the isotope enrichment at the evaporating front occurs bydiffusion through a stagnant air layer, Equation 6a withn = 1 applies. Theisotopic composition of the residual waters are characterized by their situationon a low-slope evaporation line (Figure 11). These enriched waters are then



Figure 11 Isotopic relationships during the groundwater formation in the arid and semiarid zone.


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flushed down by subsequent rains, imparting their evaporative isotope signatureto the deeper soilwaters and groundwaters.

Direct evaporation from a bare soil surface is a factor in the water balance ofan arid environment, in particular in sandy soils with a high water table [(e.g.dry lakes and diffuse groundwater discharge zones in the desert (Fontes et al1986)]. In other areas it is the transpiration flux that accounts for the bulk ofthe evapo-transpiration flux from the soil. All evidence shows that there isno fractionation between isotopes as roots take up water. On the other hand,transpiration selectively utilizes the soilwaters during the growing period inspring and summer. If, as is typically the case, the seasonal input of precipitationis not yet mixed in the soil moisture of the root zone, this results in a selectionof part of the seasonal cycle in the input to the groundwaters, resulting usuallyin more depleted isotope values in the deeper soil compared to the mean annualvalue. This property has been utilized to investigate the water balance of thegroundwater recharge process, e.g. by Simpson et al (1970).

The two processes of direct evaporation from the soil and seasonal selectionby transpiration have opposing isotopic signals. It is thus often difficult to useisotopic profiles in the soil as a measure of the evaporative history of a site.However, in combination with the salinity profile in the soil [salinity builds upas a function of the total evapo-transpiration flux (Eriksson & Khunakassem1969)] more quantitative information can be resolved.

4.5 Water in PlantsAs discussed above, soilwater is taken up by the plants essentially unfraction-ated. This feature is used extensively to characterize the source of water forthe plant and transpiration, whether soil- or riverwater (Dawson & Ehlerringer1991), snowmelt waters, or shallow or deep soilwaters in the arid zone wheremarked isotope gradients exist (Adar et al 1995). Yakir & Yehieli (1995) havealso exploited this feature to distinguish between saline groundwaters and flashfloods at the Dead Sea shore.

Leaves are the site of the evaporation front in the plant, specifically, thecells located near the stomata. These waters are highly enriched in the heavyisotopes (Gonfiantini et al 1965) when stomata are open, to an extent given bythe steady-state isotope composition with respect to the evaporation process asgiven by the formula,δSS = (1− h)δin + hδa+ ε, derived from Equation 4.Steady state can be practically achieved by virtue of the large flux/volume ratioin the transpiring leaves.

Allison et al (1985), Yakir et al (1990), and others who calculated the expectedsteady-state isotopic composition often found the overall enrichment in theleafwaters to fall short of the expected steady-state value. Allison interpretedsuch a difference to be the result of dilution by unfractionated source waters



in the veins of the leaves [amounting at times up to 30% of the total leafwatervolume (Allison et al 1985)]. In the case of cotton plants, Yakir et al (1990), onthe other hand, postulated that part of the leafwater (in the mesophyllic cells)lags behind the steady-state enrichment due to a slow water exchange with theouter water sheath, which is assumed to be in momentary steady state relativeto the evaporation process during the day and to take up unfractionated sourcewater at night.

Under steady-state conditions obviouslyδE = δin, so that the transpirationflux is unfractionated with respect to the soilwater taken up by the plants.

The enriched isotopic values of plant waters is being used to check the au-thenticity of fruit products and distinguish them from those adulterated by water(Bricout et al 1972).

5. GROUNDWATERS

5.1 Meteoric WatersThe isotopic composition of groundwaters that are recharged by direct infil-tration in the temperature zone follows that of the incident precipitation ratherclosely. The surface processes discussed above usually cause shifts in the iso-tope composition of less than 1‰ inδ(18O) (Gat & Tzur 1967). The isotopecomposition of groundwaters are thus useful tools for identifying recharge sites(utilizing geographic effects on the isotopic composition of rain, such as thealtitude effect), tracing mixing patterns and, in particular, for detecting the en-croachment of lake waters (which are enriched in the heavy isotopes) (Payne1970) of seawater and other extraneous water sources. In the arid zone, directinfiltration can result in recharge only on sand. In other cases some surfacerunoff is a necessary prerequisite for accumulating sufficient water depth toenable occurrence of recharge. Both small-scale local runoff, feeding locallyimportant aquifer pockets in limestones and crystalline rocks, and large-scaleregional flood flows, which recharge large deep regional aquifers, are importantrecharge mechanisms. Each of these recharge pathways imprints a distinguish-able stable isotopic signature on the recharge waters (Figure 11).

In order to take full advantage of the possibilities of understanding groundwaterformation, the detailed isotope effects of the recharge and runoff processes needto be established for different climate conditions and for each watershed. Suchstudies are still few and incomplete. When available, the changing isotope com-position in groundwater recharge or runoff could be used as a sensitive monitorof climatic and anthropogenic changes in the watershed (Gat & Lister 1995).

Up to temperatures of about 60–80◦C, the isotopic composition of groundwa-ters is conserved in the aquifer, so that even waters recharged in the Pleistocene


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more than 10,000 years ago are still useful for climate reconstruction of thesepast conditions based on the isotopic composition of paleowaters (Fontes 1981).For example, the wetter climate of North Africa was established by interpret-ing paleowaters in the Nubian sandstone aquifers (Sonntag et al 1976). Byvirtue of the lower d-excess value of such waters (d∼ 7‰), Merlivat & Jouzel(1979) established a higher humidity over the Atlantic at that time. However,the advantage of having a record for both oxygen and hydrogen isotopes inpaleowaters (whereas most other climate proxies only give the18O record) isoffset by the difficulty of reliably dating such waters.

5.2 Geothermal and Formation WaterAt elevated temperatures of more than 80◦C the isotopic exchange betweenthe groundwater and the host rock becomes noticeable, resulting in an “oxy-gen shift” towards more positiveδ(18O) values. However, extrapolation of theisotopic composition of such waters [at constantδ(D)] enables one, in manycases, to identify the meteoric water origin of such waters. At still higher tem-peratures and when steam venting occurs, the isotopic exchange between waterand the steam at these elevated temperatures causes changes of both oxygenand hydrogen isotopes. These relationships have been extensively utilized forthe exploration of geothermal systems and reservoirs, as reviewed by Panichi& Gonfiantini (1981).

Below the cycling meteoric waters one often encounters brackish waters andbrines in deeper geologic formations. The term “formation water” is used todenote such water that is obtained from a geologic formation and is representa-tive of its fluid content. No a priori connotation concerning the origin of salinityor of water is intended.

The isotopic composition provides information about the origin and history ofthe water phase. Recent studies have indicated that in marine sediments on thecontinents, the original marine brines have as a rule been replaced by meteoricwaters (Clayton et al 1966, Hitchon & Friedman 1969). Unfortunately, manyaspects of the mechanism of movement of water and salt remain obscure. Theevolution of such brines has been discussed by Fleisher et al (1977).

Since chloride is the dominant anion of all these brines, it is reasonable tosuppose that, initially, the salts were mainly of a marine origin.

The simplest initial solution to be considered is that of seawater or marine in-terstitial water. Another set of initial solutions that can be considered comprisescontinental surface waters. The origin of the dissolved salts in the formationwaters may then be quite varied, ranging from marine salts to residual salts ofcontinental weathering reactions. In another model, the sources of dissolvedsalts are minerals in the geologic formations, especially evaporites. In this



model, there is then no a priori relation between the dissolved salt compositionand the isotopic composition of the water.

These initial solutions are then exposed to various processes. Three geo-chemical processes can be considered. The first process is surface exposureaccompanied by evaporative water loss. According to Lloyd (1966), the iso-tope composition tends in this case towards an enriched value while salinityincreases.

A second process is that of ultrafiltration by clay membranes. This processresults in a slight enrichment of both18O and D in the residual brine, correspond-ing to a parallel increase in salinity with a preferential increase of multivalentcations relative to sodium. The single stage isotope fractionation factors indilute NaCl solutions as a result of this process are 0.9993 for18O and 0.998 fordeuterium (Coplen & Hanshaw 1973). The fractionation effect is thus smallerby one order of magnitude than the fractionation in the evaporation process.The relationships between the increase of salinity and increase in18O content byultrafiltration for the solutions considered by Coplen & Hanshaw (under condi-tions of Rayleigh fractionation) is given by the equationS= S0 exp(1δ18/0.7).If one assumes that the process starts with seawater pushed through a clay mem-brane filter until the dissolved salt increases by a factor of 4 to a concentration of150 g/liter (which are concentrations typical of formation waters), then this rela-tion predicts an enrichment of18O in the water by+ 1‰. A parallel increase inthe Ca/Na ratio would be expected (White 1965). If additional freshwater thenflows through this system, a further increase of the Ca/Na ratio may result (form-ing a CaCl2 type water), accompanied, however, by a possible slight decreasein the concentration of total dissolved solids. The18O content of the residualbrine approaches in the steady-state limit an enrichment of 0.7‰ relative to themeteoric water (this being the single stage fractionation factor for this process).

A third process that can affect the isotopic composition of formation waters isinteraction with minerals in the formation, especially at elevated temperatures.In this case the changes in the chemical and isotopic compositions depend onthe nature of the minerals. One extreme case occurs when the fluid is in contactwith limestones: The18O content of the fluid increases by exchange, withoutsignificant change of the salinity. Another extreme occurs in the presence ofanhydrous evaporite deposits, which when they dissolve, cause the salinityto increase without significantly changing the isotope composition of the fluid.White et al (1973) have claimed that in the presence of clays deuterium exchangecan occur, resulting in a slight decrease of theδ(D) values.

The three processes mentioned result in increases of salinity and enrichmentof 18O in the brine (Figure 12). Mixing with meteoric waters has the oppositeeffect, i.e. a decrease in salinity and depletion in18O. The evidence of the


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Figure 12 Isotopic changes of formation waters.

hydrogen isotopes is more ambiguous, since both exchange with clays andmixing with meteoric waters results in slightly depletedδ(D) values.

Figure 12 shows the changes in salinity and isotopic composition predictedfor the waters, as these evolve through any of the processes described above orwhen mixing with meteoric waters occurs.

6. ICE AND SNOW

Ice and snow present a special situation in isotope hydrology. Unlike the liquidraindrop, the solid phases do not exhibit isotope exchange with atmosphericmoisture but conserve the isotopic composition as formed in the cloud. Becauseof this, hailstones were used to probe the conditions within a thundercloud (Facyet al 1963, Macklin et al 1970, Jouzel et al 1975).

As shown by Jouzel & Merlivat (1984) the formation of snow shows a non-equilibrium effect that gives rise to precipitation with an elevated d-excess value.As an example, snow in Jerusalem was measured with a value of d= 32‰,compared to d' 20‰ in the ambient air. The same effect is responsible for theseasonality of d-values in precipitation of the northern latitudes, e.g. in Canada(Fritz et al 1987), with higher values in winter. It has been noted above thatthe removal of snow or hail from an air mass with which it is not in isotopicequilibrium does not satisfy the conditions for a Rayleigh process.

The accumulation of snow on the ground, with each successive snowfallkeeping its identity (with the exception of mixing caused by snowdrifts) has



enabled one to use cores in glaciers to establish the most reliable isotope recordof precipitation as far back as the Pleistocene (e.g. Dansgaard et al 1969). Theestablishment of the isotope record in glaciers constitutes one of the highlightsof isotope hydrology, which is not discussed further here in view of the manygood descriptions available.

The sublimation of the ice removes the material layer by layer. It couldbe expected then that this process would not alter the isotopic composition ofthe remaining snowpack, unlike the situation in an evaporating water body inwhich the effect of surface fractionation is propagated into the residual liquidby mixing. However, this simple model is not realized in practice, apparentlydue to diffusion of vapor through the void spaces in the snow pack and also dueto some surface melting and percolation of meltwaters into the remaining snow.As a result of such processes, which have been extensively explored (Moser &Stichler 1974, Whillans & Grootes 1985, Friedman et al 1991, also the review ofArnason 1981) meltwaters do not always show the same isotopic composition asthat of the snow; the remainder of the snow pack shows increasing enrichment ofthe heavy isotopes and reduced values of the d-parameter. An extreme exampleis reported by Claasen & Downey (1995) in the canopy-intercepted snowfallon evergreens, where enrichments of up to 2‰ inδ(18O) are found.

Another facet of solid-phase isotope hydrology is the constitution of the sea-ice cover and of lake ice. Merlivat (1978) has shown that a marine glacier ismade up mostly of accumulated precipitation and that only the very bottomlayers are frozen seawater. On a lake in northern Wisconsin, on the other hand,Bowser & Gat (1995) found the lower half of the ice cover to be ice grown inequilibrium with the lake water, which, however, is mixed with unfractionatedlake water trapped in the ice during the freezing process. The upper 50% of theice core is evidently made up of lake water that overflows on top of the subsidingice layer, mixes with the accumulated snow on the surface, and is then refrozen.

7. APPLICATIONS TO PALEOCLIMATOLOGY

The usefulness of the changes in isotope composition in the hydrologic cyclefor hydrological studies is self-evident. The essentially conservative behaviorof the isotope in the subsurface, on the one hand, and the characteristic andwell-documented changes in the atmospheric section of the cycle and at thesurface-atmosphere interface, on the other hand, provide a reliable tracer (fin-gerprint). This tracer has been utilized to determine the origin of water bodies,to quantify the water balance of surface and subsurface reservoirs, and to tracethe mechanisms of salinization of water resources (Gat 1975).

The isotope record in climatic archives is increasingly called upon to providepaleoclimatic information. This analysis is based on two premises: (a) that the


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isotope composition of the proxy material can be related to that of the water cycleof past periods and (b) that from the latter one can evaluate the relevant climaticor hydrologic parameters. Both of these premises need a critical appraisal.

As for the first of these premises, the situation is fairly straightforward wherethe proxy material is the water substance itself, as in ice cores, paleowaters, orinterstitial waters in sediment cores. However, the constancy of the isotopicrecord has to be addressed, whether with regard to diffusive mixing in thecolumn, the interaction with the rock matrix or sedimentary material (affectingthe oxygen isotopes primarily), or the effect of biogeochemical interactions thatoccur in porewaters, which affect mainly the hydrogen isotopes. The porewatersystem cannot be regarded as a closed one for more than a few thousand yearsat best because the older isotopic signals are erased by porewater diffusion (GSLister, unpublished data).

In the case of proxy materials such as lacustrine carbonates, shells of land-snails (Goodfriend et al 1989), or plant materials (cellulose), the major issueis the relationship of the isotopic composition in the proxy to the environ-mental waters. Lake carbonate deposits are related to the lake waters througha temperature-dependent fractionation factor (often compounded by an addi-tional “vital effect,” i.e. a biochemical effect). The lake waters, in turn, arerelated to the inflowing waters isotope composition and to the degree of isotopicenrichment of the lake waters.

Thus both hydrological and climatic factors are involved, as well as ecologicalones, as discussed by Talbot (1990) and Gat & Lister (1995).

Plant material, on the other hand, is related to the hydrological cycle throughthe soilwater taken up in the roots. The isotope composition is, however,modified by hydrological and physiological processes in the plant, such aswater stress, ambient humidity, and also the timing of stomata openings.

In all these cases the isotope composition of the precipitation is thus modifiedby factors that are climatically and environmentally controlled and that need tobe recognized in great detail under a variety of climatic conditions.

The second premise mentioned above—which concerns the interpretationof changes in the isotope composition of precipitation with time in terms ofchanging climate parameters, temperature in particular—is based on the strongcorrelation between the isotopic composition of precipitation and the ambienttemperature in the present-day water cycle. Various authors have attemptedto impose this relationship on the paleo-record. This obviously could be anillusive procedure when one considers that the sources and sinks of water aswell as details of the global circulation undergo important changes over theperiods concerned. The approach of modelers such as Joussaume & Jouzel(1993), who attempt to derive global isotope maps for past climate conditions



as reference points for the local climatic record, is a more promising approach.This effort, however, requires a rather detailed understanding of the bound-ary conditions both over the ocean and at the plant (soil)-atmosphere inter-face.

It is interesting to note that the interest in the paleoclimatic applications ofisotope hydrology has been a stronger motivation for the study of stable isotopesin the hydrologic cycle (yielding much useful hydrological information as a by-product) than that provided by the hydrological sciences directly.

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